2 - soil formation - mitchell soga (1)

5 CHAPTER 2 Soil Formation 2.1 INTRODUCTION The variety of geomaterials encountered in engineering problems is almost limitless, ranging from hard, dense, large pieces of rock, through gravel, sand, silt, and clay to organic deposits of soft, compressible peat. All these materials may exist over a wide range of densities and water contents. A number of different soil types may be present at any site, and the composition may vary over intervals as small as a few millimeters. It is not surprising, therefore, that much of the geoengineer’s effort is directed at the identification of soils and the evaluation of the appropriate properties for use in a particular analysis or design. Perhaps what is surprising is that the application of the principles of mechanics to a material as diverse as soil meets with as much success as it does. To understand and appreciate the characteristics of any soil deposit require an understanding of what the material is and how it reached its present state. This requires consideration of rock and soil weathering, the erosion and transportation of soil materials, deposi- tional processes, and postdepositional changes in sed- iments. Some important aspects of these processes and their effects are presented in this chapter and in Chap- ter 8. Each has been the subject of numerous books and articles, and the amount of available information is enormous. Thus, it is possible only to summarize the subject and to encourage consultation of the references for more detail. 2.2 THE EARTH’S CRUST The continental crust covers 29 percent of Earth’s sur- face. Seismic measurements indicate that the continen- tal crust is about 30 to 40 km thick, which is 6 to 8 times thicker than the crust beneath the ocean. Granitic (acid) rocks predominate beneath the continents, and basaltic (basic) rocks predominate beneath the oceans. Because of these lithologic differences, the continental crust average density of 2.7 is slightly less than the oceanic crust average density of 2.8. The elemental compositions of the whole Earth and the crust are in- dicated in Fig. 2.1. There are more than 100 elements, but 90 percent of Earth consists of iron, oxygen, sili- con, and magnesium. Less iron is found in the crust than in the core because its higher density causes it to sink. Silicon, aluminum, calcium, potassium, and so- dium are more abundant in the crust than in the core because they are lighter elements. Oxygen is the only anion that has an abundance of more than 1 percent by weight; however, it is very abundant by volume. Silicon, aluminum, magnesium, and oxygen are the most commonly observed elements in soils. Within depths up to 2 km, the rocks are 75 percent secondary (sedimentary and metamorphic) and 25 per- cent igneous. From depths of 2 to 15 km, the rocks are about 95 percent igneous and 5 percent secondary. Soils may extend from the ground surface to depths of several hundred meters. In many cases the distinction between soil and rock is difficult, as the boundary be- tween soft rock and hard soil is not precisely defined. Earth materials that fall in this range are sometimes difficult to deal with in engineering and construction, as it is not always clear whether they should be treated as soils or rocks. A temperature gradient of about 1C per 30 m exists between the bottom of Earth’s crust at 1200C and the surface. 1 The rate of cooling as molten rock magma 1 In some localized areas, usually within regions of recent crustal movement (e.g., fault lines, volcanic zones) the gradient may exceed 20C per 100 m. Such regions are of interest both because of their potential as geologic hazards and because of their possible value as sources of geothermal energy. Copyrighted Material Copyright © 2005 John Wiley & Sons Retrieved from: www.knovel.com

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Soil Formation


The variety of geomaterials encountered in engineeringproblems is almost limitless, ranging from hard, dense,large pieces of rock, through gravel, sand, silt, and clayto organic deposits of soft, compressible peat. All thesematerials may exist over a wide range of densities andwater contents. A number of different soil types maybe present at any site, and the composition may varyover intervals as small as a few millimeters.

It is not surprising, therefore, that much of thegeoengineer’s effort is directed at the identification ofsoils and the evaluation of the appropriate propertiesfor use in a particular analysis or design. Perhaps whatis surprising is that the application of the principles ofmechanics to a material as diverse as soil meets withas much success as it does.

To understand and appreciate the characteristics ofany soil deposit require an understanding of what thematerial is and how it reached its present state. Thisrequires consideration of rock and soil weathering, theerosion and transportation of soil materials, deposi-tional processes, and postdepositional changes in sed-iments. Some important aspects of these processes andtheir effects are presented in this chapter and in Chap-ter 8. Each has been the subject of numerous booksand articles, and the amount of available informationis enormous. Thus, it is possible only to summarize thesubject and to encourage consultation of the referencesfor more detail.


The continental crust covers 29 percent of Earth’s sur-face. Seismic measurements indicate that the continen-tal crust is about 30 to 40 km thick, which is 6 to 8times thicker than the crust beneath the ocean. Granitic

(acid) rocks predominate beneath the continents, andbasaltic (basic) rocks predominate beneath the oceans.Because of these lithologic differences, the continentalcrust average density of 2.7 is slightly less than theoceanic crust average density of 2.8. The elementalcompositions of the whole Earth and the crust are in-dicated in Fig. 2.1. There are more than 100 elements,but 90 percent of Earth consists of iron, oxygen, sili-con, and magnesium. Less iron is found in the crustthan in the core because its higher density causes it tosink. Silicon, aluminum, calcium, potassium, and so-dium are more abundant in the crust than in the corebecause they are lighter elements. Oxygen is the onlyanion that has an abundance of more than 1 percentby weight; however, it is very abundant by volume.Silicon, aluminum, magnesium, and oxygen are themost commonly observed elements in soils.

Within depths up to 2 km, the rocks are 75 percentsecondary (sedimentary and metamorphic) and 25 per-cent igneous. From depths of 2 to 15 km, the rocks areabout 95 percent igneous and 5 percent secondary.Soils may extend from the ground surface to depths ofseveral hundred meters. In many cases the distinctionbetween soil and rock is difficult, as the boundary be-tween soft rock and hard soil is not precisely defined.Earth materials that fall in this range are sometimesdifficult to deal with in engineering and construction,as it is not always clear whether they should be treatedas soils or rocks.

A temperature gradient of about 1�C per 30 m existsbetween the bottom of Earth’s crust at 1200�C and thesurface.1 The rate of cooling as molten rock magma

1 In some localized areas, usually within regions of recent crustalmovement (e.g., fault lines, volcanic zones) the gradient may exceed20�C per 100 m. Such regions are of interest both because of theirpotential as geologic hazards and because of their possible value assources of geothermal energy.



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Oxygen 46%

Oxygen 30%

Silicon 28%

Silicon 15%

Aluminum 8%

Aluminum 1.1%

Iron 6%

Iron 35%

Magnesium 4%Magnesium 13%

Calcium 2.4% Calcium 1.1%Potassium 2.3%Sodium 2.1%

Nickel 2.4%Sulfur 1.9%Other <1%Other <1%












Earth's Crust Whole Earth

Figure 2.1 Elemental composition of the whole Earth andthe crust (percent by weight) (data from Press and Siever,1994).

Figure 2.2 Geologic cycle.

Figure 2.3 Simplified version of the rock cycle.

moves from the interior of Earth toward the surfacehas a significant influence on the characteristics of theresulting rock. The more rapid the cooling, the smallerare the crystals that form because of the reduced timefor atoms to attain minimum energy configurations.Cooling may be so rapid in a volcanic eruption that nocrystalline structure develops before solidification, andan amorphous material such as obsidian (volcanicglass) is formed.


The surface of Earth is acted on by four basic proc-esses that proceed in a never-ending cycle, as indi-cated in Fig. 2.2. Denudation includes all of those pro-cesses that act to wear down land masses. These in-clude landslides, debris flows, avalanche transport,wind abrasion, and overland flows such as rivers andstreams. Weathering includes all of the destructive me-chanical and chemical processes that break downexisting rock masses in situ. Erosion initiates thetransportation of weathering products by variousagents from one region to another—generally fromhigh areas to low. Weathering and erosion convertrocks into sediment and form soil. Deposition involvesthe accumulation of sediments transported previously

from some other area. Sediment formation pertains toprocesses by which accumulated sediments are densi-fied, altered in composition, and converted into rock.

Crustal movement involves both gradual rising ofunloaded areas and slow subsidence of depositional ba-sins (epirogenic movements) and abrupt movements(tectonic movements) such as those associated withfaulting and earthquakes. Crustal movements may alsoresult in the formation of new rock masses throughigneous or plutonic activity. The interrelationships ofthese processes are shown in Fig. 2.3.

More than one process acts simultaneously in na-ture. For example, both weathering and erosion takeplace at the surface during periods of uplift, or oro-genic activity (mountain building), and deposition, sed-iment formation, and regional subsidence are generallycontemporaneous. This accounts in part for the widevariety of topographic and soil conditions in any area.



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2500 Precambrian



Figure 2.4 Stratigraphic timescale column. Numbers repre-sent millions of years before the present.

The stratigraphic timescale column shown in Fig.2.4 gives the sequence of rocks formed during geolog-ical time. Rocks are grouped by age into eons, eras,periods, and epochs. Each time period of the columnis represented by its appropriate system of rocks ob-served on Earth’s surface along with radioactive agedating. Among various periods, the Quaternary period(from 1.6 million years ago to the present) deservesspecial attention since the top few tens of meters ofEarth’s surface, which geotechnical engineers oftenwork in, were developed during this period. The Qua-ternary period is subdivided into the Holocene (the10,000 years after the last glacial period) and the Pleis-tocene. The deposits during this period are controlledmainly by the change in climate, as it was too short atime for any major tectonic changes to occur in thepositions of land masses and seas. There were as manyas 20 glacial and interglacial periods during the Qua-ternary. At one time, ice sheets covered more thanthree times their present extent. Worldwide sea leveloscillations due to glacial and interglacial cycles affectsoil formation (weathering, erosion, and sedimenta-tion) as well as postdepositional changes such as con-solidation and leaching.


Rocks are heterogeneous assemblages of smaller com-ponents. The smallest and chemically purest of thesecomponents are elements, which combine to form in-organic compounds of fixed composition known asminerals. Hence, rocks are composed of minerals oraggregates of minerals. Rocks are sometimes glassy(volcanic glass, obsidian, e.g.), but usually consist ofminerals that crystallized together or in sequence(metamorphic and igneous rocks), or of aggregatesof detrital components (most sedimentary rocks).Sometimes, rocks are composed entirely of one typeof mineral (say flint or rock salt), but generally theycontain many different minerals, and often the rock isa collection or aggregation of small particles that arethemselves pieces of rocks. Books on petrography maylist more than 1000 species of rock types. Fortunately,however, many of them fall into groups with similarengineering attributes, so that only about 40 rocknames will suffice for most geotechnical engineeringpurposes.

Minerals have a definite chemical composition andan ordered arrangement of components (a crystal lat-tice); a few minerals are disordered and without defin-able crystal structure (amorphous). Crystal size andstructure have an important influence on the resistanceof different rocks to weathering. Factors controlling thestability of different crystal structures are consideredin Chapter 3. The greatest electrochemical stability ofa crystal is reached at its crystallization temperature.As temperature falls below the crystallization temper-ature, the structural stability decreases. For example,olivine crystallizes from igneous rock magma at hightemperature, and it is one of the most unstable igneous-rock-forming minerals. On the other hand, quartz doesnot assume its final crystal structure until the temper-ature drops below 573�C. Because of its high stability,quartz is the most abundant nonclay mineral in soils,although it comprises only about 12 percent of igneousrocks.

As magma cools, minerals may form and remain, orthey may react progressively to form other minerals atlower temperatures. Bowen’s reaction series, shown inFig. 2.5, indicates the crystallization sequence ofthe silicate minerals as temperature decreases from1200�C. This reaction series closely parallels variousweathering stability series as shown later in Table 2.2.For example, in an intermediate granitic rock, horn-blende and plagioclase feldspar would be expected tochemically weather before orthoclase feldspar, whichwould chemically weather before muscovite mica, andso on.



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Figure 2.5 Bowen’s reaction series of mineral stability. Eachmineral is more stable than the one above it on the list.

Mineralogy textbooks commonly list determinativeproperties for about 200 minerals. The list of the mostcommon rock- or soil-forming minerals is rather short,however. Common minerals found in soils are listed inTable 2.1. The top six silicates originate from rocks byphysical weathering processes, whereas the other min-erals are formed by chemical weathering processes.Further description of important minerals found insoils is given in Chapter 3.


Weathering of rocks and soils is a destructive processwhereby debris of various sizes, compositions, andshapes is formed.2 The new compositions are usuallymore stable than the old and involve a decrease in theinternal energy of the materials. As erosion moves theground surface downward, pressures and temperaturesin the rocks are decreased, so they then possess aninternal energy above that for equilibrium in the newenvironment. This, in conjunction with exposure to theatmosphere, water, and various chemical and biologicalagents, results in processes of alteration.

A variety of physical, chemical, and biological proc-esses act to break down rock masses. Physical proc-esses reduce particle size, increase surface area, andincrease bulk volume. Chemical and biological proc-esses can cause complete changes in both physical andchemical properties.

2 A general definition of weathering (Reiche, 1945; Keller, 1957) is:the response of materials within the lithosphere to conditions at ornear its contact with the atmosphere, the hydrosphere, and perhapsmore importantly, the biosphere. The biosphere is the entire spaceoccupied by living organisms; the hydrosphere is the aqueous enve-lope of Earth; and the lithosphere is the solid part of Earth.

Physical Processes of Weathering

Physical weathering processes cause in situ breakdownwithout chemical change. Five processes are impor-tant:

1. Unloading Cracks and joints may form todepths of hundreds of meters below the groundsurface when the effective confining pressure isreduced. Reduction in confining pressure may re-sult from uplift, erosion, or changes in fluid pres-sure. Exfoliation is the spalling or peeling off ofsurface layers of rocks. Exfoliation may occurduring rock excavation and tunneling. The termpopping rock is used to describe the sudden spall-ing of rock slabs as a result of stress release.

2. Thermal Expansion and Contraction The ef-fects of thermal expansion and contraction rangefrom creation of planes of weakness from strainsalready present in a rock to complete fracture.Repeated frost and insolation (daytime heating)may be important in some desert areas. Fires cancause very rapid temperature increase and rockweathering.

3. Crystal Growth, Including Frost Action Thecrystallization pressures of salts and the pressureassociated with the freezing of water in saturatedrocks may cause significant disintegration. Manytalus deposits have been formed by frost action.However, the role of freeze–thaw in physicalweathering has been debated (Birkeland, 1984).The rapid rates and high amplitude of tempera-ture change required to produce necessary pres-sure have not been confirmed in the field. Instead,some researchers favor the process in which thinfilms of adsorbed water is the agent that promotesweathering. These films can be adsorbed sotightly that they cannot freeze. However, the wa-ter is attracted to a freezing front and pressuresexerted during the migration of these films canbreak the rock apart.

4. Colloid Plucking The shrinkage of colloidalmaterials on drying can exert a tensile stress onsurfaces with which they are in contact.3

5. Organic Activity The growth of plant roots inexisting fractures in rocks is an important weath-ering process. In addition, the activities ofworms, rodents, and humans may cause consid-erable mixing in the zone of weathering.

3 To appreciate this phenomenon, smear a film of highly plastic claypaste on the back of your hand and let it dry.



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Table 2.1 Common Soil Minerals

Name Chemical Formula Characteristics

Quartz SiO2 Abundant in sand and siltFeldspar (Na,K)AlO2[SiO2]3


Abundant in soil that is not leached extensively

Mica K2Al2O5[Si2O5]3Al4(OH)4


Source of K in most temperate-zone soils

Amphibole (Ca,Na,K)2,3(Mg,Fe,Al)5(OH)2[(Si,Al)4O11]2 Easily weathered to clay minerals and oxidesPyroxene (Ca,Mg,Fe,Ti,Al)(Si.Al)O3 Easily weatheredOlivine (Mg,Fe)2SiO4 Easily weatheredEpidoteTourmalineZirconRutileKaolinite






Highly resistant to chemical weathering; usedas ‘‘index mineral’’ in pedologic studies


Mx(Si,Al)8(Al,Fe,Mg)4O20(OH)4,where M � interlayer cation

Abundant in clays as products of weathering;source of exchangeable cations in soils

Allophane Si3Al4O12 � nH2O Abundant in soils derived from volcanic ashdeposits

Imogolite Si2Al4O10 � 5H2OGibbsite Al(OH)3 Abundant in leached soilsGoethite FeO(OH) Most abundant Fe oxideHematite Fe2O3 Abundant in warm regionFerrihydrate Fe10O15 � 9H2O Abundant in organic horizonsBirnessite (Na,Ca)Mn7O14 � 2.8H2O Most abundant Mn oxideCalcite CaCO3 Most abundant carbonateGypsum CaSO4 � 2H2O Abundant in arid regions

Adapted from Sposito (1989).

Physical weathering processes are generally theforerunners of chemical weathering. Their main con-tributions are to loosen rock masses, reduce particlesizes, and increase the available surface area for chem-ical attack.

Chemical Processes of Weathering

Chemical weathering transforms one mineral to an-other or completely dissolves the mineral. Practicallyall chemical weathering processes depend on the pres-ence of water. Hydration, that is, the surface adsorptionof water, is the forerunner of all the more complexchemical reactions, many of which proceed simulta-neously. Some important chemical processes are listedbelow.

1. Hydrolysis, probably the most important chemi-cal process, is the reaction between the mineraland H� and (OH)� of water. The small size of

the ion enables it to enter the lattice of mineralsand replace existing cations. For feldspar,Orthoclase feldspar:

� �K silicate � H OH� �→ H silicate � K OH (alkaline)


� �Ca silicate � 2H OH

→ H silicate � Ca(OH) (basic)2

As water is absorbed into feldspar, kaolinite isoften produced. In a similar way, other clay min-erals and zeolites (microporous aluminosilicates)may form by weathering of silicate minerals asthe associated ions such as silica, sodium, potas-sium, calcium, and magnesium are lost into so-



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Figure 2.6 Solubility of alumina and amorphous silica inwater (Keller, 1964b).

lution.Hydrolysis will not continue in the presence of

static water. Continued driving of the reaction tothe right requires removal of soluble materials byleaching, complexing, adsorption, and precipita-tion, as well as the continued introduction of H�

ions.Carbonic acid (H2CO3) speeds chemical

weathering. This weak acid is formed by the so-lution in rainwater of a small amount of carbondioxide gas from the atmosphere. Additional car-bonic acid and other acids are produced by theroots of plants, by insects that live in the soil,and by the bacteria that degrade plant and animalremains.

The pH of the system is important because itinfluences the amount of available H�, the solu-bility of SiO2 and Al2O3, and the type of claymineral that may form. The solubility of silicaand alumina as a function of pH is shown in Fig.2.6.

2. Chelation involves the complexing and removalof metal ions. It helps to drive hydrolysis reac-tions. For example,Muscovite:

K [Si Al ]Al O (OH) � 6C O H � 8H O2 6 2 4 20 4 2 4 2 2

� � 0 �→ 2K � 6C O Al � 6Si(OH) � 8OH2 4 4

Oxalic acid (C2O4H2), the chelating agent, re-leases C2O4

2�, which forms a soluble complexwith Al3� to enhance dissolution of muscovite.Ring-structured organic compounds derived fromhumus can act as chelating agents by holdingmetal ions within the rings by covalent bonding.

3. Cation exchange is important in chemical weath-ering in at least three ways:a. It may cause replacement of hydrogen on

hydrogen bearing colloids. This reduces theability of the colloids to bring H� to unweath-ered surfaces.

b. The ions held by Al2O3 and SiO2 colloids in-fluence the types of clay minerals that form.

c. Physical properties of the system such as thepermeability may depend on the adsorbed ionconcentrations and types.

4. Oxidation is the loss of electrons by cations, andreduction is the gain of electrons. Both are im-portant in chemical weathering. Most importantoxidation products depend on dissolved oxygenin the water. The oxidation of pyrite is typical ofmany oxidation reactions during weathering(Keller, 1957):

2FeS � 2H O � 7O → 2FeSO � 2H SO2 2 2 4 2 4

FeSO � 2H O → Fe(OH) � H SO4 2 2 2 4


Oxidation of Fe(OH)2 gives

4Fe(OH) � O � 2H O → 4Fe(OH)2 2 2 3

2Fe(OH) → Fe O � nH O (limonite)3 2 3 2

The H2SO4 formed in these reactions rejuvenatesthe process. It may also drive the hydrolysis ofsilicates and weather limestone to produce gyp-sum and carbonic acid. During the constructionof the Carsington Dam in England in the early1980s, soil in the reservoir area that containedpyrite was uncovered during construction follow-ing the excavation and exposure of air and waterof the Namurian shale used in the embankment.The sulfuric acid that was released as a result ofthe pyrite oxidation reacted with limestone toform gypsum and CO2. Accumulation of CO2 inconstruction shafts led to the asphyxiation ofworkers who were unaware of its presence. It isbelieved that the oxidation process was mediatedby bacteria (Cripps et al., 1993), as discussed fur-



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Figure 2.7 Microogranisms attached to soil particle sur-faces: (a) bacteria attached to sand particle (from Robertsonet al. 1993 in Chenu and Stotzky, 2002), (b) bacterial mi-croaggregate [from Robert and Chenu (1992) in Chenu andStotzky (2002)], and (c) biofilm on soil surface (from Chenuand Stotzky (2002).

ther in the next section.Many iron minerals weather to iron oxide

(Fe2O3, hematite). The red soils of warm, humidregions are colored by iron oxides. Oxides canact as cementing agents between soil particles.

Reduction reactions, which are of importancerelative to the influences of bacterial action andplants on weathering, store energy that may beused in later stages of weathering.

5. Carbonation is the combination of carbonate orbicarbonate ions with earth materials. Atmos-pheric CO2 is the source of the ions. Limestonemade of calcite and dolomite is one of the rocksthat weather most quickly especially in humidregions. The carbonation of dolomitic limestoneproceeds as follows:

CaMg(CO ) � 2CO � 2H O3 2 2 2

→ Ca(HCO ) � Mg(HCO )3 2 3 2

The dissolved components can be carried off inwater solution. They may also be precipitated atlocations away from the original formation.

Microbiological Effects

Several types of microorganisms are found in soils;there are cellular microorganisms (bacteria, archea, al-gae, fungi, protozoa, and slime molds) and noncellularmicroorganisms (viruses). They may be nearly round,rodlike, or spiral and range in size from less than 1 to100 �m, which is equivalent to coarse clay size to finesand size. Figure 2.7a shows bacteria adhering toquartz sand grains, and Fig. 2.7b shows clay mineralscoating around the cell envelope, forming what arecalled bacterial microaggregates.4 A few billion to 3trillion microorganisms exist in a kilogram of soil nearthe ground surface and bacteria are dominant. Micro-organisms can reproduce very rapidly. The replicationrate is controlled by factors such as temperature, pH,ionic concentrations, nutrients, and water availability.Under ideal conditions, the ‘‘generation time’’ for bac-terial fission can be as short as 10 min; however, anhour scale is typical. These high-speed generationrates, mutation, and natural selection lead to very fastadaptation and extraordinary biodiversity.

Autotrophic photosynthetic bacteria, that is, photo-autotrophs, played a crucial role in the geological de-

4 Further details of how microorganisms adhere to soil surfaces aregiven in Chenu and Stotzky (2002).



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velopment of Earth (Hattori, 1973; McCarty, 2004).Photosynthetic bacteria, cyanobacteria, or ‘‘blue-greenbacteria’’ evolved about 3.5 billion years ago (Proter-ozoic era—Precambrian), and they are the oldestknown fossils. Cyanobacteria use energy from the sunto reduce the carbon in CO2 to cellular carbon and toobtain the needed electrons for oxidizing the oxygenin water to molecular oxygen. During the Archaeanperiod (2.5 billion years ago), cyanobacteria convertedthe atmosphere from reducing to oxidizing andchanged the mineral nature of Earth.

Eukaryotic algae evolved later, followed by the mul-ticellular eukaryotes including plants. Photosynthesisis the primary producer of the organic particulate mat-ter in shale, sand, silt, and clay, as well as in coal,petroleum, and methane deposits. Furthermore, cyano-bacteria and algae increase the water pH when theyconsume CO2 dissolved in water, resulting in carbonateformation and precipitation of magnesium and calciumcarbonates, leading to Earth’s major carbonate forma-tions.

Aerobic bacteria live in the presence of dissolvedoxygen. Anaerobic bacteria survive only in the absenceof oxygen. Facultative bacteria can live with or withoutoxygen. Some bacteria may resort to fermentation tosustain their metabolism under anaerobic conditions(Purves et al., 1997). For example, in the case of an-aerobic conditions, fermenting bacteria oxidize carbo-hydrates to produce simple organic acids and H2 thatare used to reduction of ferric (Fe3�) iron, sulfate re-duction, and the generation of methane (Chapelle,2001). Microbial energy metabolism involves electrontransfers, and the electron sources and acceptors canbe both organic and inorganic compounds (Horn andMeike, 1995). Most soil bacteria derive their carbonand energy directly from organic matter and its oxi-dation. Some other bacteria derive their energy fromoxidation of inorganic substances such as ammonium,sulfur, and iron and most of their carbon from carbondioxide. Therefore, biological activity mediates geo-chemical reactions, causing them to proceed at ratesthat are sometimes orders of magnitude more rapidthan would be predicted solely on the basis of the ther-mochemical reactions involved.

Bacteria tend to adhere to mineral surfaces and formmicrocolonies known as biofilms as shown in Fig. 2.7c.Some biofilms are made of single-type bacteria, whileothers involve symbiotic communities where two ormore bacteria types coexist and complement eachother. For example, biofilms involved in rock weath-ering may involve an upper aerobic layer, followed byan intermediate facultative layer that rests on top of theaerobic layer that produces the weathering agents

(e.g., acids) directly on the rock surface (Ehrlich,1998). Biofilms bind cations in the pore fluid and fa-cilitate nucleation and crystal growth even at low ionicconcentrations in the pore fluid (Konhauser and Urru-tia, 1999). After nucleation is initiated, further mineralgrowth or precipitation can occur abiotically, includingthe precipitation of amorphous iron–aluminum sili-cates and poorly crystallized claylike minerals, such asallophone, imogolite, and smectite (Urrutia and Bev-eridge, 1995; Ehrlich, 1999; Barton et al., 2001).

In the case of the Carsington Dam construction,Cripps et al. (1993) hypothesized that autotrophic bac-teria greatly accelerated the oxidation rate of the pyrite,so that it occurred within months during construction.The resulting sulfuric acid reacted with the drainageblanket constructed of carboniferous limestone, whichthen resulted in precipitation of gypsum and iron hy-droxide, clogging of drains and generation of carbondioxide.

Weathering Products

The products of weathering, several of which will gen-erally coexist at one time, include:

1. Unaltered minerals that are either highly resistantor freshly exposed

2. Newly formed, more stable minerals having thesame structure as the original mineral

3. Newly formed minerals having a form similar tothe original, but a changed internal structure

4. Products of disrupted minerals, either at or trans-ported from the site. Such minerals might includea. Colloidal gels of Al2O3 and SiO2

b. Clay mineralsc. Zeolitesd. Cations and anions in solutione. Mineral precipitates

5. Unused guest reactants

The relationship between minerals and differentweathering stages is given in Table 2.2. The similaritybetween the order of representative minerals for thedifferent weathering stages and Bowen’s reaction se-ries given earlier (Fig. 2.5) may be noted.

Contrasts in compositions between terrestrial and lu-nar soils can be accounted for largely in terms of dif-ferences in chemical weathering. Soils on Earth arecomposed mainly of quartz and clay minerals becausethe minerals of lower stability, such as feldspar, oli-vine, hornblende, and glasses, are rapidly removed bychemical weathering. On the Moon, however, the ab-sence of water and free oxygen prevent chemicalweathering. Hence, lunar soils are made up mainly offragmented parent rock and rapidly crystallized



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Table 2.2 Representative Minerals and SoilsAssociated with Weathering Stages


RepresentativeMinerals Typical Soil Groups

Early Weathering Stages






Gypsum (also halite,sodium nitrate)

Calcite (also dolomiteapatite)

Olivine-hornblende(also pyroxenes)

Biotite (also glauco-nite, nontronite)

Albite (also anorthitemicrocline, ortho-clase)

Soils dominated bythese minerals in thefine silt and clay frac-tions are the youthfulsoils all over theworld, but mainlysoils of the desertregions where limitedwater keeps chemicalweathering to a mini-mum.

Intermediate Weathering Stages


QuartzMuscovite (also illite)2�1 layer silicates (in-

cluding vermiculite,expanded hydrousmica)


Soils dominated bythese minerals in thefine silt and clay frac-tions are mainly thoseof temperate regionsdeveloped under grassor trees. Includes themajor soils of thewheat and corn beltsof the world.

Advanced weathering stages



KaoliniteGibbsiteHematite (also geothite,

limonite)Anatase (also rutile,


Many intensely weath-ered soils of the warmand humid equatorialregions have clayfractions dominatedby these minerals.They are frequentlycharacterized by theirinfertility.

From Jackson and Sherman (1953).

glasses. Mineral fragments in lunar soils include pla-gioclase feldspar, pyroxene, ilmenite, olivine, and po-tassium feldspar. Quartz is extremely rare because it isnot abundant in the source rocks. Carrier et al. (1991)present an excellent compilation of information aboutthe composition and properties of lunar soil.

Effects of Climate, Topography, Parent Material,Time, and Biotic Factors

The rate at which weathering can proceed is controlledby parent material and climate. Topography, apart fromits influence on climate, determines primarily the rateof erosion, and this controls the depth of soil accu-mulation and the time available for weathering prior toremoval of material from the site. In areas of steeptopography, rapid mechanical weathering followed byrapid down-slope movement of the debris results information of talus slopes (piles of relatively unweath-ered coarse rock fragments).

Climate determines the amount of water present, thetemperature, and the character of the vegetative cover,and these, in turn, affect the biologic complex. Somegeneral influences of climate are:

1. For a given amount of rainfall, chemical weath-ering proceeds more rapidly in warm than in coolclimates. At normal temperatures, reaction ratesapproximately double for each 10�C rise in tem-perature.

2. At a given temperature, weathering proceedsmore rapidly in a wet climate than in a dry cli-mate provided there is good drainage.

3. The depth to the water table influences weather-ing by determining the depth to which air isavailable as a gas or in solution and by its effecton the type of biotic activity.

4. Type of rainfall is important: short, intense rainserode and run off, whereas light-intensity, long-duration rains soak in and aid in leaching.

Table 2.3 summarizes geomorphologic processes indifferent morphoclimatic zones. The nature and rate ofthese geomorphologic processes control landform as-semblages.

During the early stages of weathering and soil for-mation, the parent material is much more importantthan it is after intense weathering for long periods oftime. Climate ultimately becomes a more dominantfactor in residual soil formation than parent material.

Of the igneous rock-forming minerals, only quartzand, to a much lesser extent, feldspar, have sufficientchemical durability to persist over long periods ofweathering. Quartz is most abundant in coarse-grainedgranular rocks such as granite, granodiorite, andgneiss, where it typically occurs in grains in the mil-limeter size range. Consequently, granitic rocks are themain source of sand.

In addition to the microbiological activities dis-cussed previously, biological factors of importance in-clude the influences of vegetation on erosion rate andthe cycling of elements between plants and soils. Mi-



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Table 2.3 Morphoclimatic Zones and the Associated Geomorphologic Processes





Precipitation(mm) Relative Importance of Geomorphologic Processes

Glacial �0 0–1000 Mechanical weathering rates (especially frost action)high; chemical weathering rates low, massmovement rates low except locally; fluvial actionconfined to seasonal melt; glacial action at amaximum; wind action significant

Periglacial �1 to 2 100–1000 Mechanical weathering very active with frost action ata maximum; chemical weathering rates low tomoderate; mass movement very active; fluvialprocesses seasonally active; wind action rateslocally high. Effects of the repeated formation anddecay of permafrost.

Wet midlatitude 0–20 400–1800 Chemical weathering rates moderate, increasing tohigh at lower latitudes; mechanical weatheringactivity moderate with frost action important athigher latitudes; mass movement activity moderateto high; moderate rates of fluvial processes; windaction confined to coasts.

Dry continental 0–10 100–400 Chemical weathering rates low to moderate;mechanical weathering, especially frost action,seasonally active; mass movement moderate andepisodic; fluvial processes active in wet season;wind action locally moderate.

Hot dry (aridtropical)

10–30 0–300 Mechanical weathering rates high (especially saltweathering), chemical weathering minimum, massmovement minimal; rates of fluvial activitygenerally very low but sporadically high; windaction at maximum.

Hot semidry(semiaridtropical)

10–30 300–600 Chemical weathering rates moderate to low;mechanical weathering locally active especially ondrier and cooler margins; mass movement locallyactive but sporadic; fluvial action rates high butepisodic; wind action moderate to high.

Hot wet–dry(humid–aridtropical)

20–30 600–1500 Chemical weathering active during wet season; ratesof mechanical weathering low to moderate; massmovement fairly active; fluvial action high duringwet season with overland and channel flow; windaction generally minimum but locally moderate indry season.

Hot wet(humidtropical)

20–30 �1500 High potential rates of chemical weathering;mechanical weathering limited; active, highlyepisodic mass movement; moderate to low rates ofstream corrosion but locally high rates of dissolvedand suspended load transport.




Rates of all processes vary significantly with altitude;mechanical and glacial action becomes significant athigh elevations.

From Fookes et al. (2000).



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crobial decomposition of the heavy layers of organicmatter in top soils formed through photosynthesis re-sults in oxygen depletion and carbon oxidation back toCO2, which is leached by rainwater that penetrates intothe subsurface. The high CO2 concentration, loweredpH, and anaerobic nature of these penetrating waterscause reduction and solutioning of iron and manganeseminerals, the reduction of sulfates, and dissolution ofcarbonate rocks. If the moving waters become co-mingled with oxygenated water in the ground, or asgroundwater emerges into rivers and streams, iron,manganese, and sulfide oxidation results, and carbon-ate precipitation can occur (McCarty, 2004).

The time needed to weather different materials var-ies greatly. The more unconsolidated and permeablethe parent material, and the warmer and more humidthe climate, the shorter the time needed to achievesome given amount of soil formation. The rates ofweathering and soil development decrease with in-creasing time.

The time for soil formation from hard rock parentmaterials may be very great; however, young soils candevelop in less than 100 years from loessial, glacial,and volcanic parent material (Millar et al., 1965). Py-rite bearing rocks are known to break apart and un-dergo chemical and mineral transformations in only afew years.


There are three general mechanisms of clay formationby weathering (Eberl, 1984): (1) inheritance, (2) neo-formation, and (3) transformation. Inheritance meansthat a clay mineral originated from reactions that oc-curred in another area during a previous stage in therock cycle and that the clay is stable enough to remainin its present environment. Origin by neoformationmeans that the clay has precipitated from solution orformed from reactions of amorphous materials. Trans-formation genesis requires that the clay has kept someof its inherited structure while undergoing chemicalreactions. These reactions are typically characterizedby ion exchange with the surrounding environmentand/or layer transformation in which the structure ofoctahedral, tetrahedral, or fixed interlayer cations ismodified.

The behavior of nonclay colloids such as silica andalumina during crystallization is important in deter-mining the specific clay minerals that form. Certaingeneral principles apply.5

5 The considerations in Chapter 6 provide a basis for these statements.

1. Alkaline earths (Ca2�, Mg2�) flocculate silica.2. Alkalis (K�, Na�, Li�) disperse silica.3. Low pH flocculates colloids.4. High electrolyte content flocculates colloids.5. Aluminous suspensions are more easily floccu-

lated than siliceous suspensions.6. Dispersed phases are more easily removed by

groundwater than flocculated phases.

Factors important in determining the formation ofspecific clay minerals are discussed below. The struc-ture and detailed characterization of these minerals arecovered in Chapter 3.

Kaolinite Minerals

Kaolinite formation is favored when alumina is abun-dant and silica is scarce because of the 1�1 sil-ica�alumina structure, as opposed to the 2�1 silica toalumina structure of the three-layer minerals. Condi-tions leading to kaolinite formation usually include lowelectrolyte content, low pH, and the removal of ionsthat tend to flocculate silica (Mg, Ca, Fe) by leaching.Most kaolinite is formed from feldspars and micas byacid leaching of acidic (SiO2-rich) granitic rocks. Ka-olinite forms in areas where precipitation is relativelyhigh, and there is good drainage to ensure leaching ofcations and iron.

Halloysite forms as a result of the leaching of feld-spar by H2SO4, which is often produced by the oxi-dation of pyrite, as shown earlier. The combination ofconditions that results in halloysite formation is oftenfound in high-rain volcanic areas such as Hawaii andthe Cascade Mountains of the Pacific Northwest in theUnited States.

Smectite Minerals

Smectites, because of their 2�1 silica�alumina struc-ture, form where silica is abundant, as is the casewhere both silica and alumina are flocculated. Condi-tions favoring this are high pH, high electrolyte con-tent, and the presence of more Mg2� and Ca2� thanNa� and K�. Rocks that are high in alkaline earths,such as the basic and intermediate igneous rocks, vol-canic ash, and their derivatives containing ferromag-nesian minerals and calcic plagioclase, are usual parentmaterials. Climatic conditions where evaporation ex-ceeds precipitation and where there is poor leachingand drainage, such as in arid and semiarid areas, favorthe formation of smectite.

Illite (Hydrous Mica) and Vermiculite

Hydrous mica minerals form under conditions similarto those leading to the formation of smectites. In ad-dition, the presence of potassium is essential; so ig-



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neous or metamorphic rocks and their derivatives arethe usual parent rocks. Weathering of feldspar in coolclimates often leads to the development of illite. Al-teration of muscovite to illite and biotite to vermiculiteduring weathering is also a significant source of theseminerals. Interstratifications of vermiculite with micaand chlorite are common. The high stability of illite isresponsible for its abundance and persistence in soilsand sediments.

Chlorite Minerals

Chlorites can form by alteration of smectite throughintroduction of sufficient Mg2� to cause formation ofa brucitelike layer that replaces the interlayer water.Biotite from igneous and metamorphic rocks may alterto trioctahedral chlorites and mixed-layer chlorite–vermiculite. Chlorites also occur in low- to medium-grade metamorphic rocks and in soils derived fromsuch rocks.


The above considerations are greatly simplified, andthere are numerous ramifications, alterations, and var-iations in the processes. One clay type may transformto another by cation exchange and weathering undernew conditions. Entire structures may change, for ex-ample, from 2�1 to 1�1, so that montmorillonite formswhen magnesium-rich rocks weather under humid,moderately drained conditions, but then alters to kao-linite as leaching continues. Kaolinite does not form inthe presence of significant concentrations of calcium.

The relative proportions of potassium and magne-sium determine how much montmorillonite and illiteform. Some montmorillonites alter to illite in a marineenvironment due to the high K� concentration. Mixed-layer clays often form by partial leaching of K orMg(OH)2 from between illite and chlorite layers andby incomplete adsorption of K or Mg(OH)2 in mont-morillonite or vermiculite.

Further details of the clay minerals are given inChapter 3. More detailed discussions of clay mineralformation are given by Keller (1957, 1964a & b), Wea-ver and Pollard (1973), Eberl (1984), and Velde(1995), among others.


In situ weathering processes lead to a sequence of ho-rizons within a soil, provided erosion does not rapidlyremove soil from the site. The horizons may gradeabruptly from one to the next or be difficult to distin-

guish. Their thickness may range from a few milli-meters to several meters. The horizons may differ inany or all of the following ways:

1. Degree of breakdown of parent material2. Content and character of organic material3. Kind and amount of secondary minerals4. pH5. Particle size distribution

All the horizons considered together, including theunderlying parent material, form the soil profile.6 Thepart of the profile above the parent material is termedthe solum. Eluviation is the movement of soil materialfrom one place to another within the soil, either insolution or in suspension as a result of excess precip-itation over evaporation. Eluvial horizons have lost ma-terial; illuvial horizons have gained material.

Master horizons are designated by the capital lettersO, A, B, C, and R (Table 2.4). Subordinate symbolsare used as suffixes after the master horizon designa-tions to indicate dominant features of different kindsof horizons, as indicated in the table. The O horizonsare generally present at the soil surface under nativevegetation, but they may also be buried by sedimen-tation of alluvium, loess, or ash fall. The A horizon isthe zone of eluviation where humified organic matteraccumulates with the mineral fraction. The amount oforganic matters (fibers to humic/fulvic acids) variesfrom 0.1 percent in desert soils to 5 percent or morein organic soils and affects many engineering proper-ties including compressibility, shrinkage, strength andchemical sorption. The B horizon is the zone of illu-viation where clay, iron compounds, some resistantminerals, cations, and humus accumulate. The R ho-rizon is the consolidated rock, and the C horizon con-sists of the altered material from which A and Bhorizons are formed.

Soil profiles developed by weathering can be cate-gorized into three groups on the basis of their miner-alogy and chemical composition as shown in Fig. 2.8(Press and Siever, 1994). Pedalfers, which are formedin moist climate, are soils rich in aluminum and ironoxides and silicates such as quartz and clay minerals.All soluble minerals such as calcium carbonate isleached away. They have a thick A horizon and can befound in much of the areas of moderate to high rainfallin the eastern United States, Canada, and Europe. Ped-ocals, which are formed in dry climate, are soils rich

6 Residual soil profiles should not be confused with soil profiles re-sulting from successive deposition of strata of different soil types inalluvial, lake, or marine environments.



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Table 2.4 Designations of Master Horizons and Subordinate Symbols forHorizons of Soil Profiles

Master Horizons

O1 Organic undecomposed horizonO2 Organic decomposed horizonA1 Organic accumulation in mineral soil horizonA2 Leached bleached horizon (eluviated)A3 Transition horizon to BAB Transition horizon between A and B—more like A in upper part

A and B A2 with less than 50% of horizon occupied by spots of BAC Transition horizon, not dominated by either A or C

B and A B with less than 50% of horizon occupied by spots of A2B Horizon with accumulation of clay, iron, cations, humus; residual

concentration of clay; coatings; or alterations of originalmaterial forming clay and structure

B1 Transition horizon more like B than AB2 Maximum expression of B horizonB3 Transitional horizon to C or R

C Altered material from which A and B horizons are presumed to beformed

R Consolidated bedrock

Subordinate Symbols

b Buried horizonca Calcium in horizoncs Gypsum in horizoncn Concretions in horizon

f Frozen horizong Gleyed horizonh Humus in horizonir Iron accumulation in horizonm Cemented horizonp Plowed horizon

sa Salt accumulation in horizonsi Silica cemented horizont Clay accumulation in horizonx Fragipan horizon

II, III, IV Lithologic discontinuitiesA�2, B�2 Second sequence in bisequal soil

Adapted from Soil Survey Staff (1975).

in calcium from the calcium carbonates and other sol-uble minerals originated from sedimentary bedrock.Soil water is drawn up near the surface by evaporation,leaving calcium carbonate pellets and nodules. Theycan be found in the southwest United States. Laterite,which is formed in a wet, tropical climate, is rich inaluminum and iron oxides, iron-rich clays, and alu-minum hydroxides. Silica and calcium carbonates areleached away from the soil. It has a very thin A ho-

rizon because most of the organic matter is recycledfrom the surface to the vegetation.

Lithologic discontinuities may be common in land-scapes where erosion is severe, and these discontinui-ties are often marked by stone layers from previouserosion cycles. In some places, soils have developedseveral sequences of A and B horizons, which are su-perimposed over each other. Superimposed soil se-quences are likely the result of climate changes acting



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(a) (b) (c)




Humus andleached soil(quartz andclay mineralspresent)

Some iron andaluminium oxidesprecipitated; allsoluble materials,such ascarbonates,leached away

Granitebedrock C



Sandstone,shale, andlimestonebedrock

Calciumcarbonatepellets andnodulesprecipitated

Humus andleached soil

Thin or absenthumus

Thick masses ofinsoluble iron andaluminum oxides;occasional quartz

Iron-rich clays andaluminumhydroxides

Thin leached zone

Mafic igneousbedrock

Figure 2.8 Major soil types: (a) Pedalfer soil profile developed on granite, (b) Pedocal soilprofile developed on sedimentary bedrock, and (c) Laterite soil profile developed on maficigneous rock (from Press and Siever, 1994).

on uniform geologic materials, or are the remnants offormer soil profiles (paleosoils) that have been buriedunder younger soils (Olson, 1981).


Streams, ocean currents, waves, wind, groundwater,glaciers, and gravity continually erode and transportsoils and rock debris away from the zone of weather-ing. Each of these transporting agents may causemarked physical changes in the sediment it carries. Al-though detailed treatment of erosion, transportation,and depositional processes is outside the scope of thisbook, a brief outline of their principles and their effectson the transported soil is helpful in understanding theproperties of the transported material.


Erosion includes all processes of denudation that in-volve the wearing away of the land surface by me-

chanical action. The transporting agents, for example,water, wind, and ice, are by themselves capable onlyof limited wearing action on rocks, but the process isreinforced when these agents contain particles of thetransported material.

Transportation of sediment requires first that it bepicked up by the eroding agent. Greater average flowvelocities in the transporting medium may be requiredto erode than to transport particles. Particles are erodedwhen the drag and lift of the fluid exceed the gravi-tational, cohesive, and frictional forces acting to holdthem in place. The stream velocity required to erodedoes not decrease indefinitely with decreasing particlesize because small particles remain within the bound-ary layer adjacent to the stream bed where the actualstream velocity is much less than the average velocity.Relationships between particle size and average streamvelocity required to erode and transport particles bywind and water are shown in Fig. 2.9.

Ice has the greatest competency for sediment move-ment of all the transportation agents. There is no limitto the size of particles that may be carried. Ice pushes



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Figure 2.9 Comparison of erosion and transport curves forair and running water. The air is a slightly more effectiveerosional agent than streams for very small particles but isineffective for those larger than sand (from Garrels, 1951).

Figure 2.10 Characteristics of glaciers (from Selmer-Olsen,1964).

material along in front and erodes the bottom and sidesof the valleys through which it flows. In an active gla-cier (Fig. 2.10), there is continuous erosion and trans-port of material from the region of ice accumulationto the region of melting. A dead glacier has been cutoff from a feeding ice field.


The different agents of sediment transport are com-pared in Table 2.5. The relative effect listed in the lastcolumn of this table denotes the importance of theagent on a geological scale with respect to the overallamount of sediment moved, with one representing thegreatest amount.

Movement of sediment in suspension by wind andwater depends on the settling velocity of the particlesand the laws of fluid motion. Under laminar flow con-ditions, the settling velocity of small particles is pro-portional to the square of the particle diameter. Forlarger particles and turbulent fluid flow, the settling ve-locity is proportional to the square root of the particlediameter. Particles stay in suspension once they havebeen set in motion as long as the turbulence of thestream is greater than the settling velocity.

The largest particles that can be transported by waterare carried by traction, which consists of rolling anddragging along the boundary between the transportingagent and the ground surface. Particles intermediate insize between the suspended load and the traction loadmay be carried by saltation, in which they move by aseries of leaps and bounds. Soluble materials are car-ried in solution and may precipitate as a result ofchanged conditions. The combined effects mean thatthe concentration of sediment is not constant throughthe depth of the transporting agent but is much greaternear the stream bed than near the top. Fine particlesmay be fairly evenly distributed from top to bottom;however, coarser particles are distributed mainly withinshort distances from the bottom, as shown in Fig. 2.11,which applies to a river following a straight course.

The major effects of transportation processes on thephysical properties of sediments are sorting and ab-rasion. Sorting may be both longitudinal, which pro-duces a progressive decrease in particle size withdistance from the source as the slope flattens, and lo-cal, which produces layers or lenses with differentgrain size distributions. Reliable prediction of the sort-ing at any point along a sediment transport system iscomplicated by the fact that flow rates vary from pointto point and usually with the seasons. Consequently,very complex sequences of materials may be found inand adjacent to stream beds.

Particle size and shape may be mechanically modi-fied by abrasive processes such as grinding, impact,and crushing during transportation. The abrading ef-fects of wind are typically hundreds of times greaterthan those of water (Kuenen, 1959). In general, abra-sion changes the shape and size of gravel size particlesbut only modifies the shapes of sand and smaller sizeparticles. Water-working of sands causes rounding and



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Table 2.5 Comparison of Sediment Transport Agents

AgentType of




Eroded byAverageVelocity Areas Affected

Max Loadper m3

Type ofTransport


Streams Turbulent A few km/h Sand All land A few tens ofkilograms

Bed load,suspendedload,solution


Waves Turbulent A few km/h Sand Coastlines A few tens ofkilograms

Same asstreams


Wind Turbulent 15 km/h Sand Arid, semiarid,beaches,plowed fields

A kilogram Bed load,suspendedload


Glaciers Laminar A few m/yr Largeboulders

High latitudesand altitudes

Hundreds ofkilograms

Bed load,suspendedload,surfaceload


Groundwater Laminar A few m/yr Colloids Soluble materialand colloids

A kilogram Solution 3

Gravity cm/yr to afew m/s

Boulders Steep slopes,sensitiveclays,saturatedcohesionlesssoils,unconsolidatedrock

2000 kg Bed load 3

Adapted from Garrels (1951).

polishing of grains, and wind-driven impact can causefrosting of grains. The shape and surface character ofparticles influences a soil’s stress–deformation andstrength properties owing to their effects on packing,volume change during shear, and interparticle friction.

Basic minerals, such as the pyroxenes, amphiboles,and some feldspars, are rapidly broken down chemi-cally during transport. Quartz, which is quite stablebecause of its resistant internal structure, may be mod-ified by mechanical action, but only at a slow rate.Quartz sand grains may survive a number of successivesedimentation cycles with no more than a percent ortwo of weight loss due to abrasion.

The surface textures of quartz sand particles reflecttheir origin, as shown by the examples in Fig. 2.12 fordifferent sands, each shown to three or four magnifi-cations. The mechanical and chemical actions, associ-

ated with a beach environment, produce a relativelysmooth, pitted surface texture. Aeolian sands exhibit arougher surface texture, particularly over small dis-tances. Some, but not all, river sands may have a verysmooth particle surface that reflects the influence ofchemical action. Sand that has undergone change afterdeposition and burial is termed diagenetic sand. Itssurface texture may reflect a long and stable period ofinteraction with the groundwater. In some cases, veryrough surface textures can develop. Ottawa sand, a ma-terial that has been used for numerous geotechnicalresearch investigations, is such a material.

Some effects of transportation on sediment proper-ties are summarized in Table 2.6. The gradationalcharacteristics of sedimentary materials reflect theirtransportation mode as indicated in Fig. 2.13. Sedi-ments of different origins lie within specific zones of



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Figure 2.11 Schematic diagram of sediment concentration with depth in a transportingstream.

the figure, which are defined by the logarithm of theratio of 75 percent particle size to 25 percent particlesize and the median (50 percent) grain size.


Deposition of sediments from air and water is con-trolled by the same laws as their transportation. If thestream velocity and turbulence fall below the valuesneeded to keep particles in suspension or moving withthe bed load, then the particles will settle. When icemelts, the sediments may be deposited in place or car-ried away by meltwater. Materials in solution canprecipitate when exposed to conditions of changedtemperature or chemical composition, or as a result ofevaporation of water. Sediments may be divided intothose formed primarily by chemical and biologicalmeans and those composed primarily of mineral androck fragments. The latter are sometimes referred to asdetrital or clastic deposits.

The deposition of sediments into most areas is cy-clical. Some causes of cyclic deposition are:

1. Periodic earth movements2. Climatic cycles of various lengths, most notably

the annual rhythm3. Cyclic shifting of tributaries on a delta4. Periodic volcanism

The thickness of deposits formed during any onecycle may vary from less than a millimeter to hundredsof meters. The period may range from months tothousands of years, and only one or many types ofsediments may be involved.

One of the best known sediments formed by cyclicaldeposition is varved clay. Varved clays formed in gla-cial lakes during the ice retreat stage. Each layer con-sists of a lighter-colored, summer-deposited clayey siltgrading into a darker winter-deposited silty clay.Spring and summer thaws contributed clay and silt-laden meltwater to the lake. The coarsest particles set-tled first to form the summer layer. Because of themuch slower settling velocity of the clay particles,most did not settle out until the quiet winter period. Aphotograph of a vertical section through a varved clayis shown in Fig. 2.14. The alternating coarser-grained,light-colored layers and finer-grained, darker layers areclearly visible. The shear resistance along horizontalvarves is much less than that across the varves. Also,the hydraulic conductivity is much greater in the hor-izontal direction than in the vertical direction. Exten-sive deposits of varved clays are found in the northeastand north central United States and eastern Canada.Detailed description of the geology and engineeringproperties of Connecticut Valley varved clay is givenby DeGroot and Lutenegger (2003).

Complex soil deposition processes occur alongcoastlines, estuaries, and shallow shelves in relation to



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Figure 2.12 Surface textures of four sands of differing origins: (a) river sand, (b) beachsand, (c) aeolian sand and (d) diagenetic sand (courtesy of Norris, 1975).

the location of the shoreline. Soil deposits include fore-shore sand and gravels, which are sorted by wave ac-tions, organic deposits, and clays preserved in lagoons,offshore fine sands, and muds. River channels may beoverdeepened, and soft sediments then accumulate toform buried valleys. Most coastlines and estuaries ofthe world were subject to sea level changes in the Qua-ternary period. In particular, the post glacial rise of sealevel, which ended about 6000 years ago, has had aworldwide influence on the present-day coastal forms.Figure 2.15 shows alternating layers of marine (Ma)and fluvial (Diluvial-D) sediments in the geotechnicalprofile down to 400 m depth below sea level at OsakaBay, Japan (Tanaka and Locat, 1999). The observedvariation corresponds well to the local relative sea levelduring its geological history up to 1 million years ago.

Chemical and biochemical sediments may consist ofone or two kinds of materials. For example, calciumcarbonate sediments are made of calcite, which origi-nates from the shells of organisms in the deep sea (Fig.2.16a). Some clays contain significant amounts of mi-crofossils due to the depositional environment asshown in Fig. 2.16b; such clays include Mexico Cityclay (Diaz-Rodriguez et al., 1998), Ariake clay (Oht-subo et al., 1995), and Osaka Bay clay (Tanaka andLocat, 1999). The microfossils include diatoms (sili-ceous skeleton of eukarya cells in either freshwater ormarine environments), radiolaria (found in marine en-vironments and consisting mostly of silica), and for-manifera (calcium carbonate shell secreted by marineeukarya). The presence of microfossils can have a pro-found effect on the behavior of the soil mass, confer-



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Figure 2.12 (Continued )

ring unusual geotechnical properties that deviate fromgeneral property expectations, including high porosity,high liquid limit, unusual compressibility, and uniquelyhigh friction angle. For examples, see Tanaka and Lo-cat (1999) and Locat and Tanaka (2001).

While streams and rivers produce deposits accordingto grain size, a glacier transports the finest dust andlarge boulders side by side at the same rate of move-ment. If the material remains unsorted after deposition,it is called till. A mixture of all grain sizes from boul-ders to clays is known as boulder clay, which is adifficult material to work with because large bouldersmay damage excavation equipment.

Loess, which is a nonstratified aeolian deposit, isprobably the single most abundant Quaternary depositon land. It consists of silt with some small fraction ofclay, sand, and carbonate. It originated during the Qua-

ternary period from glacial out wash and deglaciatedtill areas. The deposits are spread widely and blanketpreexisting landforms. The deposits are up to 30 mthick in the Missouri and Rhine River Valleys, morethan 180 m thick in Tajikistan, and up to 330 m thickin northern China.

Depositional Environment

The environment of deposition determines the complexof physical, chemical, and biological conditions underwhich sediments accumulate and consolidate. Thethree general geographical depositional environmentsare continental, mixed continental and marine, and ma-rine. Continental deposits are located above the tidalreach and include terrestrial, paludal (swamp), andlacustrine (lake) sediments. Mixed continental and



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Table 2.6 Effects of Transportation on Sediments

Water Air Ice Gravity

Size Reduction through solution, littleabrasion in suspended load, someabrasion and impact in tractionload


Considerablegrinding andimpact


Shape androundness

Rounding of sand and gravel High degree ofrounding

Angular, soledparticles

Angular, non-spherical

Surface texture Sand: smooth, polished, shinySilt: little effect

Impact producesfrostedsurfaces

Striated surfaces Striated surfaces

Sorting Considerable sorting Very considerablesorting(progressive)

Very little sorting No sorting

Adapted from Lambe and Whitman (1969).

Figure 2.13 Influence of geologic history on sorting of particle sizes (adapted from Selmer-Olsen, 1964).



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Figure 2.14 Vertical section through varved clay from theNew Jersey meadowlands (courtesy of S. Saxena).

Figure 2.15 Soil profile of Osaka Bay showing alternatingmarine (Ma) and fluvial (Diluvial-D) layers (modified fromTanaka and Locat, 1999).

Figure 2.16 Biochemical sediments: (a) Dogs Bay calciumcarbonate sand (courtesy of E. T. Bowman) and (b) diatomsobserved in Osaka Bay clay (courtesy of Y. Watabe).

marine deposits include littoral (between the tides),deltaic, and estuarine sediments. Marine deposits arelocated below the tidal reach and consist of continentalshelf (neritic), continental slope and rise (bathyal), anddeep ocean (abyssal) sediments. Table 2.7 summarizesmain soil deposits that are formed in various types of

environments (Locat et al., 2003). Characteristic soiltypes and properties associated with these depositionalenvironments are described in Chapter 8.


Between the time a sediment is first laid down and thetime it is encountered in connection with some humanactivity, it may have been altered as a result of theaction of any one or more of several postdepositionalprocesses. These processes can be physical, chemical,and/or biological. They occur because the young sed-iment is not necessarily stable in its new environment



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Table 2.7 Depositional Environment of Various Soil Deposits

Deposits Environment Type Texture

Transported Air Aeolian sand SandWaterShallow river Fluvial (glacio-) Sand and gravel

Alluvial (glacio-) Silt and sandShallow lake Littoral Sand and gravel

Muskeg Peat—organicDeep lake Lacustrine (glacio-) Silt and clay

Flow deposits Clay to gravelMarls Silt (fossils)

Shallow ocean Estuarine Silt and clayLittoral Silt and sandShelf Silt and clay

Deep ocean Pelagic Silt and clayOozes—calcareous Silt and clayOozes—siliceous Silt and clayFlow Clay to gravel

Glacier Subglacial till Clay to bouldersSupraglacial till Sand to boulders

Residual Land Tropical soils Clay to sandSaprolite Clay to bouldersDecomposed granite Clay to bouldersColluvial soils Clay to boulders

Chemical and biochemical Lake Evaporites (sakkas)Sea Evaporites

LimestoneGas hydrates

Adapted from Locat et al. (2003).

where the material is exposed to new conditions oftemperature, pressure, and chemistry. An understand-ing of postdepositional changes is essential for under-standing of properties, interpreting soil profile data,and in reconstructing geologic history. A brief outlineof the processes is presented here; their effects on en-gineering properties are described in more detail inChapter 8.


The drying of fine-grained sediments is usually accom-panied by shrinkage and cracking. Precompression ofthe upper portions of clay layers by drying is fre-quently observed. The effects of desiccation on thestrength and water content variations with depth inLondon clay from the Thames estuary are shown inFig. 2.17. Care must be exercised in interpreting pro-files of this type because drying is only one of severalpossible causes of apparent overconsolidation (precon-

solidation pressure greater than present overburden ef-fective pressure) at shallow depths. Other importantmechanisms include partial consolidation under in-creased overburden and the effects of weathering.


Weathering and soil-forming processes are initiated innew sedimentary deposits after exposure to the atmo-sphere, just as they are on freshly exposed rock. Insome instances, weathering can result in improvementin properties or protection of underlying material. Forexample, the weathering of uplifted marine clays canlead to the replacement of sodium by potassium as thedominant exchange cation (Moum and Rosenqvist,1957). This increases both the undisturbed and re-molded strength. Water content and strength data for aNorwegian marine clay profile are shown in Fig. 2.18.It may be seen that the upper 5 m of clay, which havebeen weathered, have water content and strength var-



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Figure 2.17 Properties of Thames estuary clay. The overconsolidation in the upper 10 ftwas caused by surface drying (Skempton and Henkel, 1953).



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Figure 2.18 Clay characteristics at Manglerud in Oslo, Norway (Bjerrum, 1954).

iation characteristics similar to those of the Thamesestuary clay (see Fig. 2.17). In the case of the Nor-wegian clay, however, the plasticity values have alsochanged in the upper 5 m, providing evidence ofchanged composition. Weathering of the surface ofsome loess deposits has resulted in the formation of arelatively impervious loam that protects the underlyingmetastable loess structure from the deleterious effectsof water.

Consolidation and Densification

Consolidation (termed compaction in geology) of fine-grained sediments occurs from increased overburden,drying, or changes in the groundwater level so that theeffective stress on the material is increased. Depositsof granular material may be affected to some extent inthe same way. More significant densification of cohe-sionless soil occurs, however, as a result of dynamicloading such as induced by earthquakes or the activi-



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ties of humans. The usual effects of consolidation areto increase strength, decrease compressibility, increaseswell potential, and decrease permeability.

Even under constant effective stress conditions,structural readjustments and small compressions maycontinue for long periods owing to the viscous natureof soil structures. This ‘‘secondary compression’’ pro-vides an additional source of increased strength withtime.


Erosion of overlying sediments due to glacial processleads to mechanical overconsolidation. A typical ex-ample of this is London clay, a marine clay depositedduring the Eocene period. The erosion took place inlate Tertiary and Pleistocene times and the amount oferosion is estimated to be about 150 m in Essex(Skempton, 1961) to 300 m in the Wraybury district(Bishop et al., 1965). After the unloading, small re-loading occurred by new deposition of gravels in thelate Quaternary period. Within the London clay, fivemajor transgressive–regressive cycles are recognizedduring its deposition. The postdepositional processesare site specific; that is, the degree of weathering anddesiccation and the amount of erosion vary dependingon location. This variation in depositional and post-depositional processes results in complex mechanicalbehavior (Hight et al., 2003).

Authigenesis, Diagenesis, Cementation, andRecrystallization

Authigenesis is the formation of new minerals in placeafter deposition. Authigenesis can make grains moreangular, lower the void ratio, and decrease the per-meability. Small crystals and rock fragments may growinto aggregates of coarser particles.

Diagenesis refers to such phenomena as changes inparticle surface texture, the conversion of mineralsfrom one type to another, and the formation of inter-particle bonds as a result of increased temperature,pressure, and time. Many diagenetic changes are con-trolled by the pH and redox potential of the deposi-tional environment. With increasing depth of burial ina sedimentary basin, clayey sediments may undergosubstantial transformation. Expansive clay mineralscan transform to a nonexpansive form, for example,montmorillonite to mixed layer to illite, as a result ofthe progressive removal of water layers under pressure(Burst, 1969). Burial depths of 1000 to 5000 m maybe required, and the transformation process appearsthermally activated as a result of the increased tem-perature at these depths. Chlorite can form in mud andshale during deep burial (Weaver and Pollard, 1973).

The long-term stability of different clay minerals underconditions of elevated temperature and pressure and indifferent chemical environments is important relativeto the use of clays as containment barriers for nuclearand toxic wastes. Diagenesis studies of locked sandsshow crystal overgrowths caused by pressure solutionand compaction (Barton, 1993; Richards and Burton,1999).

Cementation has important effects on the propertiesand stability of many soil materials. Cementation is notalways easily identified, nor are its effects always read-ily determined quantitatively. It is known to contributeto clay sensitivity, and it may be responsible for anapparent preconsolidation pressure. Removal of ironcompounds from a very sensitive clay from Labrador,Canada, by leaching led to a 30-ton/m2 decrease inapparent preconsolidation pressure (Kenney et al.,1967). Coop and Airey (2003) show for carbonate soilsthat cementation develops soon after deposition andenables the soil to maintain a loose structure.

Failure to recognize cementation has resulted in con-struction disputes. For example, a soil on a major proj-ect was marked on the contract drawings as glacialtill. It proved to be so hard that it had to be blasted.The contractor claimed the soil was cemented becauseduring digging failure took place through pebbles aswell as the clay matrix. The owner concluded that thishappened because the pebbles were weathered. Properevaluation of the material before the award of the con-tract could have avoided the problem.

Clay particles adhere to the surfaces of larger siltand sand particles, a process called clay bounding.Eventually the larger grains become embedded into aclay matrix and their influence on the geotechnical be-havior becomes limited. The clay bounding providesarching of interparticle forces, maintaining a large voidratio even at high effective stresses.

Time Effects

Even freshly deposited or densified sands can developsignificant increases in strength and stiffness over rel-atively short time periods, that is, by a factor of 2 ormore within a few months (Mitchell and Solymar,1984). Time effects and the aging of both cohesiveand cohesionless soils are analyzed and reviewedby Schmertmann (1991). Uncertainty remains as towhether the mechanisms for the observed increases inapparent preconsolidation pressure, strength, and stiff-ness are chemical, physical, or both. Research is con-tinuing on this important aspect of soil behavior so thatit will be possible to predict both the amount and therate of property changes for use in the analysis of geo-technical problems. The aging process is of particular



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interest in connection with hydraulic fills and groundimprovement projects, more details are given in Chap-ter 12.

Leaching, Ion Exchange, and Differential Solution

Postdepositional changes in pore fluid chemistry canresult from the percolation of different fluids througha deposit. This may change the forces between col-loidal particles. For example, the uplift of marine clayabove sea level followed by freshwater leaching canlead to both ion exchange and the removal of dissolvedsalts. This process is important in the formation ofhighly sensitive, quick clays, as discussed in more de-tail in Chapter 8.

Materials can be removed from sediments by differ-ential solution and subsequent leaching. Calcareousand gypsiferous sediments are particularly susceptibleto solution, resulting in the formation of channels, sinkholes, and cavities.

Jointing and Fissuring of Clay Soils

Some normally consolidated clays, almost all flood-plain clays, and many preconsolidated clays are weak-ened by joints and fissures. Joints in floodplain claysresult from deposition followed by cyclic expansionand contraction from wetting and drying. Joints andfissures in preconsolidated clays result from unloadingor from shrinkage cracks during drying. Closelyspaced joints in these types of clays may contribute toslides some years after excavation of cuts. The unload-ing enables joints to open, water to enter, and the clayto soften.

Fissures have been found in normally consolidatedclays at high water contents that could not have beencaused by drying or unloading (Skempton and Northey,1952), and increased brittleness has been observed insoft clay chunks that have been stored for some time.These effects may be caused by syneresis, which is themutual attraction of clay particles to form closely knitaggregates with fissures between them. Similar behav-ior is many times observed in gelatin after aging.Weathering and the release of potassium may also re-sult in fissuring.

Vegetation, especially large trees, can cause shrink-age and fissuring of clays (Barber, 1958; Holtz, 1983).The root systems suck up water, causing large capillaryshrinkage pressures. When rain falls on crusted surfacelayers of dried-up saline lakes, it is rapidly absorbedby capillarity. The air may become so compressed thatit causes tension cracking or blowouts in a form similarin appearance to root holes. These sediments may alsoundergo severe cracking, apparently as a result of

shock. Cracks up to 2 ft wide, of unknown depth, andspaced several meters apart have caused damage tobuildings and highways.

Biological Effects

Biological activity affects soil particles by modifyingtheir arrangement, aggregating them, weathering min-eral surfaces, mediating oxidation–reduction reactions,contributing to precipitation and dissolution of miner-als, and degrading organic particles. The survival andactivity of microorganisms are controlled partly bypore geometry and local physicochemical conditions.Therefore, apart from its impact on life itself, biolog-ical activity has influenced the evolution of the earthsurface, impacted mineral, sediment, and rock forma-tion, accelerated the rate of rock weathering and al-tered its products, influenced the composition ofgroundwater, and participated in the formation of gasand petroleum hydrocarbons.

Bioturbance refers to the action of organisms livingon or in sediments. By organic cementation, they mod-ify grain size, density, or cohesion (Richardson et al.,1985; Locat et al., 2003). The aggregation activity ofvarious worms densifies deposits by changing the grainsize of the sediment. Tubes that form can provide localdrainage and decrease the bulk density. The active zoneof bioturbance is usually to depths less than 30 cm.Sticky organic mucus or polymer bridging binds to-gether clay–silt particles, producing clusters.

Chemical transformation processes are mediated byorganisms. Some notable processes are summarized asfollows (Mitchell and Santamarina, 2005):

1. Sulfur Cycle Elemental sulfur (S0) and sulfides(S2�) are the stable forms of sulfur under anaer-obic conditions, whereas sulfates (SO4

2�) are thestable forms of sulfur under aerobic conditions.Sulfides form under anaerobic conditions fromsulfates already present in seawater and sedi-ments or introduced by diffusion and ground-water flow. The sulfate ion is not reduced tosulfide at Earth surface temperature and pressureunless biologically mediated. Sulfate-reducingbacteria are anaerobic and grow best at neutralpH but are known to exist over a broad range ofpH and salt content. When exposed to aerobicconditions, reduced sulfur compounds, hydrogensulfides (H2S), and elemental sulfur are used asan energy source by sulfide-oxidizing bacteriaand converted to sulfates.

2. Iron Cycle Iron in the subsurface exists pre-dominantly in the reduced or ferrous (Fe2�) state



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or the oxidized ferric (Fe3�) state. Several micro-organisms such as the genus Thiobacillus medi-ate the iron oxidation reaction. Chapelle (2001)notes that bacteria are able to derive only relativelittle energy from oxidizing Fe2�; therefore, theymust process large amounts of Fe2� and producelarge amounts of Fe3� to obtain sufficient energyto sustain their growth.

One important consequence of the rapid oxidationof iron sulfide in the presence of oxygen is the for-mation of acid rock drainage. Although Fe(OH)3 haslow solubility, the formation of H2SO4 provides asource of important reactions in the solid and pore wa-ter phases. The total dissolved solids increases owingto the dissolution of carbonates in the soil. Gypsumcan form, with an associated volume increase, at theexpense of carbonate minerals. The precipitated ferrichydroxide is thermodynamically unstable and rapidlytransforms to yellow goethite, FeO–OH. Geothite,while stable under wet conditions, will slowly dehy-drate to red hematite, Fe2O3, under dry conditions.

Microorganisms have a limited effect on the for-mation of coarse grains. However, bioactivity can af-fect diagenetic evolution, promote the precipitation ofcementing agents, cause internal weathering, and alterfines migration, filter performance, and drainage insilts and sands.

Severely water-limited environments distress micro-organisms and hinder biological activity. Nonetheless,there is great bacterial activity in the unsaturated or-ganic surface layer of a soil where plant roots arefound. Fierer et al. (2002) observed that bacterial ac-tivity decreases by 1 or 2 orders of magnitude by 2 mof depth. Horn and Meike (1995) conclude that micro-bial activity requires 60 to 80 percent saturation.Hence, there is less reduction in bacterial count withdepth in saturated sediments. Hindered biological ac-tivities in unsaturated soils may reflect lack of nutrientsin isolated water at menisci, slow nutrient flow in per-colating water paths, and increased ionic concentrationin the pore fluid as water evaporates and dissolved saltsapproach ion saturation conditions.

The physical scales over which the physicochemical,bioorganic, and burial diagenetic processes act rangefrom atomic dimensions to kilometers, and the time-scales range from microseconds to years. Table 2.8summarizes the processes, fabric characteristics, andscales associated with different mechanisms.

Human Effects

The global human population has grown from approx-imately 600 million at the beginning of the eighteenth

century to close to 6 billion today. Human activitiesare now at such a scale as to rival forces of nature intheir influence on soil changes. The activities includerapid changes in land use and the associated landforms,soil erosion related to forest removal, and soil contam-ination by urbanization, mining, and agricultural activ-ities. Ten to 15 percent of Earth’s land surface isoccupied by industrial areas and agriculture, and an-other 6 to 8 percent is pasture land (Vitousek et al.,1997).

Mine wastes are the largest waste volumes producedby humankind. On October 21, 1966, 144 people, 116of them children, were killed when a tip of coal wasteslid onto the village of Aberfan in South Wales, UnitedKingdom. The collapse was caused by tipping of coalwaste over a natural underground spring, and the coalslag slowly turned into a liquid slurry. The tragedy wascaused by two days of continual heavy rain looseningthe coal slag. As a result of the disaster at Aberfan,the Mines and Quarries Tips Act of 1969 was intro-duced. This act was passed in order to prevent disusedtips from becoming a danger to members of the public.

Over 8000 million tons of ore have been mined inthe South African deep-level underground gold miningindustry (Blight et al., 2000). Considerations for dis-posing these wastes into tailings ponds and dams in-clude the physicochemical nature of the extractedminerals as well as the topography and climate of thedisposal sites. Tailings dams have failed, resulting indestructive mudflows (Blight, 1997). One reported casewas the failure of the Merriespruit ring-dyke gold tail-ings dam in South Africa in 1994, which killed 17people in a village nearby. Overtopping of the tailingsdyke occurred after a significant rainfall event, and ap-proximately 500,000 m3 of tailings flowed through thisbreach. The liquefied tailings flowed for a distance ofabout 2 km. A large volume of tailings was in a me-tastable state in situ, and overtopping and erosion ofthe impoundment wall exposed this material, resultingin static liquefaction of the tailings and a consequentflow failure (Fourie et al., 2001).

The urban underground in major cities is congestedby utility lines, tunnels, and building foundations.Much may be more than 100 years old; for example,more than 50 percent of the water supply pipes in Lon-don were built using cast-iron during Victorian time.Aging infrastructure changes the in situ stress condi-tion, as well as groundwater chemistry, and this canlead to changes in the stress–strain–time behavior ofthe subsoil. Underground openings are sources or sinksof different environments; tunnels can act as a ground-water drain as well as source for air into the ground.



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Table 2.8 Summary of Processes and of the Fabric Signature and Temporal Scales Associated with VariousMechanisms

Processes Mechanisms


(predominant) ScalesPhysical

Time Remarks

Physicochemical Electromechanical EF Atomic andmolecular to� 4 �m

�s to ms Two particles may rotateFF

Thermomechanical FF(some EF)

Molecular to �

0.2 mmms to min Initial contacts EF then

rotations to FF:common in selectiveenvironments

Interface dynamics FF and EF �m to �� 0.5mm

s Some large compoundparticles may bepossible at highconcentrations

Bioorganic Biomechanical EF � 0.5 mm to� 2.0 mm

s to min Some FF possibleduring bioturbation

Biophysical EE and FF �m to mm s to min Some very large clayorganic complexespossible

Biochemical Nonunique(unknown)

�m to mm h to yr New chemicals formed,some altered


Mass gravity FF localizedswirl

cm to km � yr Can operate over largephysical scales



Molecular � yr New minerals formed,some altered, changesin morphology

aEF, edge-to-face; EE, edge-to-edge; FF, face-to-face.Adapted from Mitchell and Santamarina (2005) and Bennett et al. (1991).

Detailed studies of the geotechnical impacts of suchproblems have, so far, been limited (e.g., Gourvenecet al., 2005), and further studies of the impacts of agingon existing infrastructure are needed.


Knowledge of geologic and soil-forming processesaids in anticipating and understanding the probablecomposition, structure, properties, and behavior of asoil. Along with site investigation data, characteri-zation of the landforms, that is, understanding of theformer and current geomorphological processes asso-ciated with the past and present climatic conditions,often helps to define ground conditions for designinggeotechnical structures and anticipating the long-termperformance. For example, the knowledge can be used

to infer clay mineral types, to detect the presence oforganic and high clay content layers, to locate borrowmaterials for construction, and to estimate the depth tounaltered parent material. Pedological data can be usedto surmise compositions and soil physical properties.

Transported soils are sorted, abraded, and have par-ticle surface textures that reflect the transporting me-dium. Conditions of sedimentation and the depositionalenvironment influence the grain size, size distribution,and grain arrangement. Thus, knowledge of the trans-portation and deposition history provides insight intogeotechnical engineering properties.

In short, the soil and its properties with which wedeal today are a direct and predictable consequence ofthe parent material of many years ago and of all thethings that have happened to it since. The better ourknowledge of what that parent material was and whatthe intervening events have been, the better our ability



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to deal with the soil as an engineering material. Severalexamples are given in this chapter and more are givenin Chapter 8.


1. At what depth below the ground surface does quartzstart to crystallize?

2. What are some likely consequences of the differentphysical and chemical weathering processes on themechanical and flow properties of the rocks andsoils on which they act?

3. Describe the chemical reactions of pyrite oxidationand explain how bacteria can mediate the chemicalprocesses.

4. Discuss what types of clay minerals are likely to beproduced under each morphoclimatic zone listed inTable 2.3.

5. Using Stokes’s law, derive the sedimentation speedsof spherical particles with different sizes in fresh-water under hydrostatic condition. Would theychange in saltwater? Compare the results to the datagiven in Fig. 2.9 and discuss the comparison.

6. List and discuss human activities that may poten-tially change the properties of soils.

7. Compare and contrast soil-forming processes onEarth and on the Moon in terms of the compositionand engineering properties of the soils. Explainsimilarities and differences. What is the relative im-portance of physical, chemical, and biological soil-forming processes on the Moon and on Earth?Why?

8. Considering rock and mineral stability, the typesand characteristics of weathering processes, and theimpacts of weathering on properties, what types ofearth materials would you consider most suitablefor use as chemical, radioactive, and mixed (chem-ical and radioactive) waste containment barriers?Why?

9. Prepare diagrams showing your estimates as a func-tion of elevation of the following soil characteristicsthat you would expect to encounter between thebottom and the top of Mount Kilimanjaro in Tan-zania. Give a brief explanation for each.a. Soil plasticityb. Soil gradation and mean particle sizec. Angularity–roundness of sand and gravel parti-

clesd. Iron contente. Cementation between particlesf. Organic matter contentg. Water content



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